Surface circulation patterns and the pathways of sea surface carbon dioxide (CO<sub>2</sub>) off northern Chile (~27.5° S) between 30 and 10 kyr BP: global and/or local forcing?

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CPD

6, 347–368, 2010

Pathways of sea surface paleo-CO₂ off

northern Chile

J. A. Placencia et al.

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Clim. Past Discuss., 6, 347–368, 2010 www.clim-past-discuss.net/6/347/2010/

Climate of the Past Discussions

This discussion paper is/has been under review for the journal Climate of the Past (CP). Please refer to the corresponding final paper in CP if available.

Surface circulation patterns and the

pathways of sea surface carbon dioxide

(CO

₂

) o

_ff

northern Chile (

_∼

27.5

◦

S)

between 30 and 10 kyr BP: global and/or

local forcing?

J. A. Placencia1,2,3, N. Harada4, R. Torres5, C. B. Lange2, and D. Hebbeln6

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Interactive Discussion 2

Department of Oceanography and Center for Oceanographic Research in the eastern South Pacific (FONDAP-COPAS), University of Concepci ´on, Concepci ´on, Chile

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Department of Environmental Chemistry, Faculty of Sciences, Universidad Cat ólica de la Sant´ısima Concepci ón, Concepci ón, Chile

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Research Institute for Global Change, Japan Agency for Marine-Earth Science and Technology (JAMSTEC), Yokosuka, Japan

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Centro de Investigaci ´on en Ecosistemas de la Patagonia (CIEP), Universidad Austral de Chile, Coyhaique, Chile

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MARUM Center for Marine Environmental Sciences, Bremen University, Bremen, Germany

Received: 25 February 2010 – Accepted: 1 March 2010 – Published: 31 March 2010

Correspondence to: J. A. Placencia ([email protected])

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Interactive Discussion Abstract

We present a reconstruction of past changes in partial pressure of CO2 (pCO2) from

northern Chile (∼27◦S), between 10 and 30 kyr BP, based on carbon isotope

com-position of C_37:2-alkenone. The high-pCO₂ during the entire time series indicates that northern Chile upwelling system has been a permanent source of CO2 to the

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atmosphere. The multiproxy reconstruction suggests that the CO2outgassing and

se-questration pathways were modulated by local and global mechanisms. During global glacial conditions, an enhanced coastal upwelling forcing resulted in high-availability of deep water macronutrients and a CO2-supersaturated water column, which combined

with high-inputs of iron from the continent, intensified the carbon sequestration

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way of the biological pump, through diatom biomass export. During the deglacial, a decrease in the upwelling forcing, an increment in water column stability and reduced continental inputs of iron are consistent with a larger role of calcifying organisms in the plankton assemblage in terms of carbon sequestration pathway through the carbonate system.

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1 Introduction

Coastal upwelling areas are often described as intense sources of CO2and

macronu-trients (e.g., Borges et al., 2005). The CO2-supersaturated and high-macronutrient

concentration of upwelled waters stimulate intense blooms of phytoplankton that se-quester inorganic carbon thus reducing the outgassing of CO₂ (e.g., Simpson and

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Zirino, 1980).

Large parts of the ocean have been described as High-Nutrient/Low-Chlorophyll (HNLC) areas, where phytoplankton growth is limited by the availability of iron (e.g., Martin, 1990; de Baar and Boyd, 2000; Bruland et al., 2005). Experimental evidence suggests that the availability of dissolved iron permits the rapid interception of upwelled

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and Ampuero, 2009). High-levels of dissolved iron in coastal upwelled waters are deter-mined by the input of iron into the photic layer from other sources (e.g., dust deposition, river discharges, continental shelf) in addition to the initial dissolved iron contained in the water masses that feed the upwelling (Johnson et al., 1999; Bruland et al., 2005).

Upwelling favorable winds are permanent at ∼27◦S offChile during all year

(Shaf-5

fer et al., 1999). In this area of narrow continental shelf, the recently upwelled water is characterized by low-temperature, high-nutrient concentrations, low-chlorophyll and supersaturation in CO₂(e.g., Torres et al., 1999; Torres and Ampuero, 2009). At these windy latitudes (27◦–30◦S) (Shaffer et al., 1999), the supersaturation in CO2allows for

strong-CO2fluxes from the ocean to the atmosphere (e.g., Torres et al., 2002). Vertical

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fluxes of organic matter and water column stratification are key factors that explain the intensity/extension of aerobic and anaerobic production of CO2. The intense

denitri-fication reported for upwelling areas off Peru and Chile highlights an additional input of CO₂associated with anaerobic respiration of organic matter (e.g., Friederich et al., 2008; Paulmier et al., 2008).

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Paleoceanographic reconstructions off northern Chile described a strong preces-sional cyclicity in the pattern of iron flux, suggesting that iron availability is a key factor for past changes in productivity and sequestration of atmospheric CO2 (Lamy et al.,

1998; Dezileau et al., 2004). This was related to increased precipitation during pre-cessional maxima, higher river runofftransporting iron-rich terrigenous material to the

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ocean inducing high-biological productivity (i.e., diatom blooms) and consequently, a rapid decrease in sea surface waterpCO2 (Dezileau et al., 2004). In contrast,

low-precipitation during precessional minima decreased the iron supply to the ocean, thus limiting the primary productivity and enhancing the supersaturation of CO2in the

sur-face water column (Dezileau et al., 2004).

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This study is a first approximation of millennialpCO₂changes in the eastern South Pacific (ESP) margin between 30 and 10 kyr BP, based on carbon isotope composi-tion of C37:2-alkenone. Our results reveal that important changes in the marine CO2

outgassing were modulated by local and global mechanisms, including water column stability, iron-limitation, and consequently the productivity patterns, which potentially

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control the different pathways of removal of the marine CO2from surface waters to the

deep-ocean.

2 Material and methods

We used deep-sea sediment core GeoB 3375-1 (27◦28′S, 71◦15′W; 1947 m water depth; core length 489 cm), collected from the Chilean margin. Sediment samples

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were sub-sampled at 5 cm, freeze-dried and extracted for alkenone analyses according to the methodology described by Prahl et al. (1988) and Harada et al. (2003).

1–2 grams of freeze-dried sediment was ground into a powder with an agate mortar and pestle. The bulk organic fractions were extracted from sediment samples with an accelerated solvent extractor (ASE-200, Dionex Japan, Ltd.) at 100◦C and 1000 psi

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using a mixture of dichloromethane and methanol (99:1 v/v) as the solvent. The yield of the extraction procedure was>95% for a ketone standard (2-nonadecanone).

The extracts were saponified in 0.5 M KOH in methanol at 80◦C for 2 h. The neutral fraction was recovered with a pipette, dissolved in hexane, and then separated into subfractions by silica gel column chromatography using an automatic solid preparation

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system (Rapid Trace SPE Workstation, Zymark, UK). Standards for the subfractiona-tion were hydrocarbons (C20–C38 hydrocarbons and 5α-cholestane, GLSciences Inc.,

Japan), a ketone (2-nonadecanone, Fluka Chemie, Switzerland), and a sterol (choles-terol, GLSciences Inc., Japan). The solvents used were 4 mL of hexane for fraction 1; a mixture of 2 mL of a hexane-toluene mixture (3:1 v/v), 2 mL of hexane-toluene (1:1 v/v),

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acetate (4:1 v/v) for fraction 3; and 2 mL of hexane-ethyl acetate (4:1 v/v) followed by 2 mL of ethyl acetate for fraction 4. The alkenones eluted in fraction 3. The recovery of a 2-nonadecanone standard with this procedure was>90%. A sample of the alkenone fraction was analyzed by capillary gas chromatography with a Fisons 8000 series gas chromatograph (CE Instruments, USA) equipped with a 50 m long, 0.32 mm internal

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diameter Chrompack CP-Sil 5CB fused silica column, a cold on column injector, and a flame ionization detector. The temperature regime for gas chromatography was as follows: held at 60◦C for 1 min, increased to 200◦C at 20◦C min−1, increased to 305◦C at 15◦C min−1, held at 305◦C for 32 min, increased to 320◦C at 15◦C min−1, and held at 320◦C for 5 min.

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U₃₇K′-index calculations were conducted according to Prahl and Wakeham (1987), and U₃₇K′ values were converted to alkenone-based temperature (U₃₇K′-SST) using the equation of Prahl et al. (1988), i.e.,U₃₇K′=0.034T+0.039, whereT=temperature (◦_C).

The C37:2-alkenone isotope analyses were performed using a HP6890 gas

chro-matography unit and a Finnigan MAT 252 (Thermo Electron Co. Bremen, Germany)

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mass spectrometer unit. The13C/12C isotopic ratio was also measured in an-C₃₆ hy-drocarbon (GLSciences Inc., Japan) and National Bureau of Standards standard 19 (International Atomic Energy Agency) as an internal standard. The analytical accuracy ofδ13C_C37:2-alkenone was±0.05‰.

Alkenone-CO2estimates were based on the expression:

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εp32:7=εf−b/[CO2aq] (1)

whereεp32:7 values are based on the isotopic difference between an estimate of the

δ13C values of the organic matter produced by ancient algae (δ13Corg) and the

co-existing dissolved CO2 in the photic zone (δ 13

CCO2aq) (Freeman and Pagani, 2005).

δ13C_org is determined by adding 4.2‰ to the measured δ13C_37:2 values, as a

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gare et al., 1997; Popp et al., 1998; Pagani et al., 1999). δ13CCO2aq is reconstructed

from the carbon isotope composition of planktic foraminifera, which assumes that iso-topic and chemical equilibrium prevailed among all the carbon aqueous species and atmospheric CO2 as well as foraminiferal calcite (e.g., Pagani et al., 1999). We have

selected the median of the estimate for ε_f (26‰) with an uncertainty of 1‰, derived

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from chemostat culture experiments (e.g., Laws et al., 1995; Popp et al., 1998). The b value represents an integration of all the physiological variables that affect the total carbon isotope fractionation during photosynthesis (e.g., growth rate and cell geome-try (Jasper et al., 1994; Bidigare et al., 1997). In this study, the b-value was calculated using the geometric mean regression of all available data (Bidigare et al., 1997, 1999;

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Eek et al., 1999; Popp et al., 1999; Laws et al., 2001) and mean annual phosphate concentration from the NOAA World Ocean Atlas 2001 (Conkright et al., 2002) for the core-site location (27.5◦S).

Values of CO2aq were converted to pCO2 by applying Henry’s Law (Weiss, 1974),

calculated assuming an annual mean salinity of 34.6 psu (Conkright et al., 2002) for

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the core-site (27.5◦S), and alkenone-based temperature (Lyle et al., 1992).

The age model was based on seven14C AMS dates and linear interpolation between the age control points in the upper half of the core (from 40 to 9 kyr BP) (see Mohtadi and Hebbeln, 2004).

3 Results and discussion

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3.1 Sea surface temperature reconstruction and possible oceanographic

scenarios

At 27.5◦_{S o}

ffChile, theU₃₇K′-SST of core GeoB 3375-1 exhibits a millennial-scale vari-ation of the order of ∼4◦C during the last 30 kyr BP (Fig. 1a). The core highlights

an overall warming trend from 17◦C to 20◦C between ∼30 and 17 kyr BP (Fig. 1a).

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latter was a very abrupt cooling of almost 4◦_{C which was followed by a marked}

in-crease in SST during the deglacial. Alkenone-SST and theδ18O-N. pachyderma (sin.) records in core GeoB 3375-1 are similar between 30 and 20 kyr BP in terms of general pattern and cold events, and between 15 and 10 kyr BP with respect to the deglacial warming (Fig. 1a, b). They differ markedly in the interval 20–15 kyr BP: the

alkenone-5

SST warming trend (of∼1.5◦C) around 19–17 kyr BP and the large and abrupt 16–

17 kyr BP cooling are not seen in theδ18O record. This di_fference between surface waters (recorded by alkenones which are biosynthesized by some algae of the class Haptophyceae; e.g., Prahl et al., 1988) and deeper waters (recorded byN. pachyderma (sin.) assumed to calcify at depth below the pycnocline; e.g. Andersson et al., 2009)

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may point to enhanced stratification of the water column during this time. This was previously proposed by De Pol-Holz et al. (2006) who invoked a collapse of thermo-cline ventilation due to reduced subduction rates of Eastern South Pacific Intermediate Water (ESPIW) triggered by the melting of the Patagonian Ice Sheet, promoting nitro-gen removal that potentially contributed to the changes in atmospheric CO2during the

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deglacial (De Pol-Holz et al., 2007).

The deglacial warming (between ∼15 and 10 kyr BP, Fig. 1a and b) indicates

sub-stantial changes in subsurface water conditions in the ESP with an increased south-ward advection of the tropical-sourced subsurface waters, suggesting a reorganization of the system that lags the onset of the initial SST warming recorded in core GeoB3375

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(Fig. 1a) and other cores offnorthern and central Chile (Mohtadi and Hebbeln, 2004). The reorganization included reestablishing upwelling and ventilation, with the conse-quence of a mixed and CO2-supersaturated water column (Fig. 1c).

On the other hand, the Last Glacial Maximum (LGM) is not observed in the alkenone-SST and δ18O records. Previous studies suggested that during the coldest events

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the formation rates of ESPIW and ventilation of the shallow oceanic layer offnorthern Chile (De Pol-Holz et al., 2006). On this basis, the enhanced upwelling and ventilation during the LGM would be reflected in a CO2-supersaturated water column (Fig. 1c, see

also discussion next section).

3.2 The pathways of sea surface carbon dioxide (pCO₂)

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The measured C_37:2-alkenone δ13C values ranged from −29.5‰ to −20.5‰, and

the resulting isotopic fractionation (εp) and sea surface pCO2 ranged from 5.7‰

to 15.9‰ (Table 1S, data supplement: http://www.clim-past-discuss.net/6/347/2010/ cpd-6-347-2010-supplement.pdf) and from 200.7 µatm to 415.7 µatm, respectively, for the timeframe 30 to 10 kyr BP (Fig. 1c). The sea surfacepCO2showed a notable

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coupling with the atmospheric CO2 record (e.g., Inderm ¨uhle et al., 2000) during the

last glacial and the early deglacial (Fig. 1c and d). The sea surfacepCO₂ displayed a gradual increase between ∼30 and 28 kyr BP (from 250 to 350 µatm), followed by

a gradual decrease between∼27 and 24.5 kyr BP (from 346 to 235 µatm). Between

24 and∼17 kyr BP, while atmospheric CO2 was relatively low and stable (∼200 ppmv,

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Fig. 1d), highest sea surface pCO2 values (350–400 µatm, Fig. 1c) suggest that the

water column remained supersaturated-CO2until stratified conditions stopped

ventila-tion of the water column (De Pol-Holz et al., 2006). Consequently, thepCO₂gradually declined and this lasted until ∼14 kyr BP (Fig. 1c). In contrast, atmospheric CO2

in-creased steadily from∼18 kyr BP to the start of the Holocene (Fig. 1d).

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In summary, the sea surface pCO₂ exhibited higher levels than atmospheric CO₂, highlighting that the upwelling system offnorthern Chile (27.5◦S) may have been con-tinuously supersaturated with CO2(Fig. 1c and d), and thus would have been an

unin-terrupted source of CO₂to the atmosphere during the period 30–10 kyr BP.

The sea surface pCO2 is affected by changes in temperature, salinity, total CO2

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by biological processes (i.e., photosynthesis, respiration and calcification) and by up-welling of subsurface waters enriched in respired CO₂ and nutrients (e.g., Zeebe and Wolf-Gladrow, 2001; Anderson et al., 2009).

Given that the changes in the sea surface oceanpCO2levels could result in

impor-tant changes in calcareous and/or organic carbon production rates (due to changes

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in DIC availability; e.g., Zeebe and Wolf-Gladrow, 2001), we analyzed the potential uptake pathways of marine CO2between 30 and 10 kyr BP offnorthern Chile.

We interpret the record of δ13C of planktonic foraminifera N. pachyderma (sin.) (Fig. 1e) in terms of DIC availability (e.g., K ¨ohler et al., 2005) for carbonaceous orga-nisms instead of using it solely as an upwelling strength proxy, as suggested for this

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region by Dezileau et al. (2004). Theδ13C ofN. pachyderma (sin.)record displayed a similar pattern to the carbonate (CaCO3) fluxes (Fig. 1e and f). In general terms, when

δ13C is enriched in 13C, the carbonate rain rate decreases. The relative abundance ofG. bulloides(Fig. 1g) also exhibited a similar pattern since the LGM and especially since∼15 kyr BP during the reorganization period when the three variables highlight

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the highest carbonate production, in contrast to the late glacial (Fig. 1e, f and g). We consider that the glacial increase in organic carbon export as well as biogenic opal fluxes (Fig. 1i and j) occurs under iron-enriched conditions (Fig. 1h). The response of inorganic/organic to the sea surface CO2 concentrations (Fig. 2) suggests that

or-ganic carbon production was favored during this time. On this basis, the iron availability

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is a critical factor that determines the diatom biological pump as a pathway of carbon (Fig. 1c) to the deeper waters in the form of diatom organic carbon (Fig. 1i and j). In fact, a local experiment developed in the coastal upwelling areas offChile evidenced that diatom growth and their capacity to lower the high-pCO2of recently-upwelled

wa-ters, depend on additional inputs of iron (Torres and Ampuero, 2009).

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A large reduction inpCO₂occurred between∼27 and 24 kyr BP, during the gradual

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diatoms, to rapidly lower thepCO₂ surface levels, activating a strong biological pump able to rapidly reduce the intense CO2outgassing. Furthermore, the impoverishment

in13C of theδ13C-N. pachyderma (sin.) (Fig. 1e), indicates high-DIC availability, which was destined mainly to the photosynthesis process. Moreover, the response of inor-ganic/organic carbon to CO₂ concentration (Fig. 2) describes that whereas CO₂

dis-5

solution decreases in the sea surface waters, the organic carbon production remained above carbonate production.

Between ∼24 and 17 kyr BP the water column remained highly supersaturated in

CO2 without significant depletion, due to more pCO2 released to the atmosphere

through the upwelling process. During this period, the high-export of iron-rich

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genous material (Dezileau et al., 2004) permitted an increase of the total productivity, reflected in the high-organic carbon (Fig. 1i), -biogenic opal (Fig. 1j) and -carbonate (Fig. 1f) fluxes. However, the high-total production in the water column was not enough to counteract the high-concentrations of dissolved CO₂in the water.

During the deglacial warming (∼17–16 kyr BP, see gray hachured bar in Fig. 1), the

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well-stratified water column and the collapse of the ESPIW formation (De Pol-Holz et al., 2006) apparently favored the decrease inpCO₂(Fig. 1c). We expected that surface stratification not only would help maintain primary producers close to the surface (by limiting the depth of turbulent mixing) but also isolate the mixed layer from CO2-rich

and nutrients-rich water located below the pycnocline. In fact, a slight reduction in

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organic carbon and opal fluxes is also observed during this period (Fig. 1i and j). On the other hand, all productivity proxies for calcareous organisms displayed a slight increase during this period (Fig. 1f–g), probably favoring a reduction in water column DIC (Fig. 1e).

A simultaneous intensification of the upwelling in the Southern Ocean (SO) would

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plain the deglacial rise in the atmospheric CO₂(Fig. 1d) and the high-inputs of nutrients to the SO (Anderson et al., 2009). We would expect that the intense upwelling and high-opal fluxes recorded in the∼16–10 kyr BP period in the SO (Anderson et al., 2009),

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Subantarctic Mode Water, SAMW). However, the strong-iron limitation associated to deep upwelled waters implies that the waters formed in this period were impoverished in iron and silicates, but enriched in nitrate (Anderson et al., 2009). Therefore, a ten-dency to develop a secondary silicate limitation (Hutchins and Brunland, 1998; Takeda, 1998; Martin-J ´ez ´equel et al., 2000) while still retaining high-nitrate (Sarmiento et al.,

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2007). Thus, the upwelling of SAMW (originally formed from upper Circumpolar Deep Water) by wind driven upwelling systems is expected to lead silicate limitation, and con-sequently non-siliceous functional groups dominated the phytoplankton assemblage during this period, increasing their influence in the biological pump. This scenario is supported by the decoupling recorded between organic carbon (Fig. 1i), total nitrogen

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(Fig. 1j) and opal fluxes (Fig. 1j), during∼16–10 kyr BP. Certainly, the maximum export

of organic carbon during this period (∼16–10 kyr BP) corresponds to both a maximum

of CaCO3(Fig. 1f) and minimum of opal fluxes (Fig. 1j).

The contrast between the trends of biogenic opal and organic carbon fluxes between

∼16 and 10 kyr BP off northern Chile could be reconciled by invoking a shift in the

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phytoplankton communities from diatomaceous to carbonate production. This would have reduced and increased relative contributions of organic carbon and carbonate to total export production, respectively (Fig. 1i and j).

The apparent contrast (i.e., late glacial/deglacial) becomes clear if we consider how the decrease in terrigenous iron inputs (Fig. 1h) a_ffected the nitrate, as well as silicic

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acid utilization by diatoms, reducing the total nitrogen and opal fluxes (Fig. 1j), and consequently their role in the marine CO2drawdown by the biological pump (Fig. 1i).

High-carbonate drawdown (Fig. 1f and g) at warm-SST (Fig. 1a) and high-atmospheric CO2 (Fig. 1d), suggest that the marine DIC was captured intensively by

the carbonate system, enhancing the carbonate precipitation of the calcifying

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nisms in the plankton assemblage (Fig. 2). On basis of experimental results from Iglesias-Rodriguez et al. (2008), probably, the deglacial rise of the atmospheric CO2

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(sin.) signal (Fig. 1e) confirms this hypothesis, since the elevated impoverishment in

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C could be indicative of a major and rapid consumption of DIC by calcifying orga-nisms.

The calcification increment reduces the efficiency of the marine ecosystem to store sea surface CO₂, due to changes in the bu_ffering capacity of the ocean’s carbonate

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system, as well as to draw down the atmospheric CO2 (Cerme ˜no et al., 2008). In

fact, if (1) calcium carbonate precipitation increases the marine pCO2 (e.g., Zeebe

and Wolf-Gladrow, 2001) and (2) a high-pCO₂ increases the calcification rates (e.g. Iglesias-Rodriguez et al., 2008), a positive feedback will tend to increase the calcifica-tion processes. On this basis, only biologically active iron-enrichment condicalcifica-tions would

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enhance the competitive advantages of diatoms to break the positive feedback and shift the system to capture the marine CO2pathway biological pump.

Our findings suggest that the coastal ocean off northern Chile has been a perma-nent and strong source of CO₂ to the atmosphere between 30 and 10 kyr BP. This CO2 outgassing has been modulated by local (upwelling strength) and global

mecha-15

nisms (e.g., ocean circulation, atmospheric CO2) that have changed the water column

properties (e.g., water column stability, macro- and micro-nutrients availability), and consequently driven the role and importance of the biological pump and the marine carbonate system in the sequestration of inorganic carbon to the deep-ocean.

4 Concluding remarks

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The sea surface pCO2 reconstruction indicates that the coastal ocean off northern

Chile (∼27.5◦S) has been a permanent and strong upwelling source of CO2to the

at-mosphere between 30 and 10 kyr BP. The multiproxy reconstruction reveals that the CO2 outgassing and sequestration was modulated by local and global mechanisms.

During the global glacial period, the water column remained highly supersaturated in

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the carbon sequestration pathway biological pump through diatom biomass export. In contrast, during the deglacial warming the iron-limitation altered significantly the marine CO2 sequestration pathway biological pump. Warm-SST and high-atmospheric CO2

conditions suggest that the carbon was captured intensively by the carbonate system, enhancing the carbonate precipitation of the calcifying organisms in the plankton

as-5

semblage. Overall findings suggest that biologically active iron-availability conditions controlled the competitive advantages of plankton to shift the system to capture the marine CO₂ pathway biological pump and/or carbonate system off northern Chile in the eastern south Pacific margin between 30 and 10 kyr BP.

Acknowledgements. We thank to M. Sato (JAMSTEC) for her valuable help with alkenone 10

analyses. This work was funded by the JAMSTEC Institution and the FONDAP-COPAS Center (Grant No. 150100007). J.A. Placencia was supported by scholarships granted by the Ministry of Education-Chile (Grant MECESUP/UCO-0002), the Graduate School at University of Con-cepcion, and by the POGO-IOC-SCOR fellowship 2005.

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Sea Res., Part I, 49, 1413–1429, 2002.

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Fig. 1. Core GeoB3375-1 off northern Chile. A multiproxy reconstruction between 30 and 10 kyr BP:(a)Alkenone-based SST (this study). (b)δ18O of N. pachyderma (sin)(Dezileau et al., 2004). (c) Sea surfacepCO₂ (this study). (d)Atmospheric CO₂ concentration in the Antarctic Taylor Dome ice core (composite data from Smith et al., 1999, and Inderm ¨uhle et al., 2000). (e) δ13C ofN. pachyderma (sin) (Dezileau et al., 2004). (f) *CaCO₃ fluxes. (g) The relative abundance ofG. bulloides(Mohtadi and Hebbeln, 2004). (h)*Iron fluxes (black doted line) and the Fe/Al ratio (gray line). The Fe/Al ratio indicates marine vs. terrigenous sources of iron (Dezileau et al., 2004). (i)*Organic carbon fluxes. (j)*Biogenic Opal and *Total Nitrogen fluxes. Gray hatchured bar indicates the onset of increased stratification off 30◦_{S (De}

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Surface circulation patterns and the pathways of sea surface carbon dioxide (CO<sub>2</sub>) off northern Chile (~27.5&deg; S) between 30 and 10 kyr BP: global and/or local forcing?

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Surface circulation patterns and the

pathways of sea surface carbon dioxide

(CO

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between 30 and 10 kyr BP: global and/or

local forcing?

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Surface circulation patterns and the pathways of sea surface carbon dioxide (CO<sub>2</sub>) off northern Chile (~27.5° S) between 30 and 10 kyr BP: global and/or local forcing?

₂

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